The Earth’s tungsten budget during mantle melting and crust formation

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Geochimica et Cosmochimica Acta 75 (2011) 2119–2136 www.elsevier.com/locate/gca

The Earth’s tungsten budget during mantle melting and crust formation S. Ko¨nig a,b,⇑, C. Mu¨nker b, S. Hohl a, H. Paulick a, A.R. Barth a,b, M. Lagos c, J. Pfa¨nder d, A. Bu¨chl e,1 a

Rheinische Friedrich–Wilhelms-Universita¨t Bonn, Steinmann Institut fu¨r Geologie, Mineralogie und Pala¨ontologie, Poppelsdorfer Schloss, 53115 Bonn, Germany b Universita¨t zu Ko¨ln, Institut fu¨r Geologie und Mineralogie, Zu¨lpicher Str. 49b, 50674 Ko¨ln, Germany c Institut fu¨r Nukleare Entsorgung, Karlsruher Institut fu¨r Technologie, Campus Nord, Hermann–von-Helmholtz-Platz 1, 76344 Eggenstein-Leopoldshafen, Germany d Technische Universita¨t Bergakademie Freiberg, Institut fu¨r Geologie, Bernhard–von-Cotta Str. 2, 09599 Freiberg, Germany e Max-Planck-Institut fu¨r Chemie, Abteilung Geochemie, Postfach 3060, 55020 Mainz, Germany Received 25 June 2010; accepted in revised form 20 January 2011; available online 28 January 2011

Abstract During silicate melting on Earth, W is one of the most incompatible trace elements, similar to Th, Ba or U. As W is also moderately siderophile during metal segregation, ratios of W and the lithophile Th and U in silicate rocks have therefore been used to constrain the W abundance of the Earth’s mantle and the Hf–W age of core formation. This study presents highprecision W concentration data obtained by isotope dilution for samples covering important silicate reservoirs on Earth. The data reveal significant fractionations of W from other highly incompatible lithophile elements such as Th, U, and Ta. Many arc lavas exhibit a selective enrichment of W relative to Th, U, and Nb–Ta, reflecting W enrichment in the sub-arc mantle via fluid-like components derived from subducting plates. In contrast, during enrichment by melt-like subduction components, W is generally slightly depleted relative to Th and U, but is still enriched relative to Ta. Hence, all arc rocks and the continental crust exhibit uniformly low Ta/W (ca. 1), whereas W/Th and W/U may show opposite fractionation trends, depending on the role of fluid- and melt-like subduction components. Further high-precision W data for OIBs and MORBs reveal a systematic depletion of W in both rock types relative to other HFSE, resulting in high Ta/W that are complementary to the low Ta/W observed in arc rocks and the continental crust. Similar to previous interpretations based on Nb/U and Ce/Pb systematics, our Ta/W data confirm a depletion of the depleted upper mantle (DM) in fluid mobile elements relative to the primitive mantle (PRIMA). The abundance of W in the depleted upper mantle relative to other immobile and highly incompatible elements such as Nb and Ta is therefore not representative of the bulk silicate Earth. Based on mass balance calculations using Ta–W systematics in the major silicate reservoirs, the W abundance of the Earth’s primitive mantle can be constrained to 12 ppb, resulting in revised ratios of W–U and W–Th of 0.53 and 0.14, respectively. The newly constrained Hf–W ratio of the silicate Earth is 25.8, significantly higher than previously estimated (18.7) and overlaps within error the Hf–W ratio proposed for the Moon (ca. 24.9). The 182Hf–182W model age for the formation of the Earth’s core that is inferred from the 182W abundance and the Hf/W of the silicate Earth is therefore younger than previously calculated, by up to 5 Myrs after solar system formation depending on the accretion models used. The similar Hf/W ratios and 182W compositions of the Earth and the silicate

⇑ Corresponding author at: Rheinische Friedrich–Wilhelms-Universita¨t Bonn, Steinmann Institut fu¨r Geologie, Mineralogie und

Pala¨ontologie, Poppelsdorfer Schloss, 53115 Bonn, Germany. E-mail address: [email protected] (S. Ko¨nig). 1 Present address: Physikalisches Institut – Didaktik der Physik, Staudtstraße 7, 91058 Erlangen, Germany. 0016-7037/$ - see front matter Ó 2011 Elsevier Ltd. All rights reserved. doi:10.1016/j.gca.2011.01.031

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Moon suggest a strong link between the Moon forming giant impact and final metal–silicate equilibration on the Earth. Ó 2011 Elsevier Ltd. All rights reserved.

1. INTRODUCTION During differentiation of the Earth into metal core and silicate mantle, W behaves as a moderately siderophile element. Hence, most of the W is sequestered into the Earth’s core (>90%, Jagoutz et al., 1979; Sun, 1982; Newsom and Palme, 1984) with much lower abundances in the primitive mantle (Newsom et al., 1996). The segregation of the Earth’s metallic core from the silicate mantle occurred probably within the first 70 million years after formation of the solar system (e.g., Yin et al., 2002; Halliday, 2004; Kleine et al., 2004; Jacobsen, 2005 and references therein). After core formation and in the absence of a residual metallic phase, W is known to behave as a highly incompatible lithophile trace element during mantle melting and silicate differentiation, similar to Ba, Th, or U (e.g., Rammensee and Wa¨nke, 1977; Newsom and Drake, 1979; Palme and Rammensee, 1981; Newsom et al., 1996). Previously reported W–Th–U data for oceanic basalts appear to indicate a similar compatibility of these three elements during mantle melting (Newsom et al., 1996). The pioneering study by Newsom et al. (1996) proposed that W/Th ratios of important silicate reservoirs on the Earth are nearly constant, suggesting a nearly similar incompatibility of both elements. In a more recent study, Arevalo and McDonough (2008) contend that W is similarly compatible than U, but more compatible than Ba and Th. Assuming near constant W/U (Arevalo and McDonough, 2008) or W/Th in mafic rocks (Newsom et al., 1996), primitive mantle values for W/U of 0.65 and W/Th of 0.15–0.19 have been employed to estimate a range of W concentration of the Earth’s primitive mantle from 13 ppb (Arevalo and McDonough, 2008) to 16 ppb (Newsom et al., 1996). This estimate of the W concentration in the primitive mantle is also the basis for the Hf–W model age for the formation of the Earth’s core that is inferred from the 182W abundance and the Hf/W of the silicate Earth (e.g., Kleine et al., 2002; Scho¨nberg et al., 2002; Yin et al., 2002). The Hf/W of the silicate Earth has been derived assuming near constant W/Th and chondritic Hf/Th ratios (Newsom et al., 1996; Kleine et al., 2002; Scho¨nberg et al., 2002; Yin et al., 2002). In contrast to previous findings (Newsom et al., 1996; Noll et al., 1996; Arevalo and McDonough, 2008), a high-precision W study on arc lavas (Ko¨nig et al., 2008) has proposed the selective enrichment of W relative to Th, U, and Ta during subduction zone processes involving fluid-like subduction components derived from subducted oceanic crust. This observation is in accord with a previous study on weathered alluvial sediments suggesting a “hydrolithophile” behaviour of W (Kamber et al., 2005). As subduction-zone magmatism significantly contributes to the growth of continental crust, continued crustal growth may have potentially led to systematically higher W/Th, W/U and lower Ta/W in the continental crust. Complementary to the growing continental crust, the upper mantle might have progressively

been depleted in W over geological time. Thus, the W/Th, W/U and Ta/W composition of the present day upper mantle may not be representative of the primitive mantle that is not affected by subduction zone processes and melt extraction (e.g., Hofmann, 1986). Following the model proposed by Ko¨nig et al. (2008), the W abundance of the primitive mantle might be substantially lower and the Hf/W of the bulk silicate Earth higher than previously estimated. In order to re-assess the budget of W in the silicate Earth, we analysed a suite of well characterized mafic rocks from a variety of tectonic settings for their W concentrations, employing isotope dilution and MC–ICPMS. The study also continues a global survey of HFSE concentrations in MORB (Bu¨chl et al., 2002; Mu¨nker et al., 2003) and OIB (Pfa¨nder et al., 2007). On the basis of our new highprecision data, the behaviour of W relative to Ba–Th–U and other HFSE during mantle melting is evaluated. As a consequence, the W budget of the bulk silicate Earth can be assessed more accurately, and its bearing on the Hf–W age of the Earth’s core is re-evaluated. 2. SAMPLE SELECTION AND ANALYTICAL TECHNIQUES The sample suite analysed in this study comprises a representative range of ocean island basalts (OIB, n = 23), mid-ocean-ridge basalts (MORB, n = 41), and subductionrelated volcanic rocks that originate from Cenozoic arc settings (herein referred to as arc rocks, n = 111). All samples (i.e., unaltered glasses and lavas) are mafic in composition (4–15 wt% MgO). The OIB suite comprises HIMU, FOZO and enriched mantle (EM II) endmembers as defined by Zindler and Hart (1986) and Hofmann (1997). HIMU-type samples are from Rurutu (young and old volcanic lavas) and Tubuai Island in the Cook-Austral chain (Chauvel et al., 1992, 1995). EM II-type samples are from Maupiti and Raiatea in the Society chain (Blais et al., 2002) and from Savai’i and Upolu Island of western Samoa (Workman et al., 2004). Major, trace element, and Sr, Nd, Pb isotope compositions are given by Chauvel et al. (1995) for Rurutu, by Chauvel et al. (1992) for Tubuai, by Workman et al. (2004) for Samoa, and by Blais et al. (2002) for the Society Islands. A complete dataset including high-precision (isotope dilution) high field strength element data (Zr, Hf, Nb, Ta) as well as Hf isotope compositions for these OIB samples is provided by Pfa¨nder et al. (2007). High precision high field strength element data for the 7.30°S to 11°30S MAR glasses (MAR South) are presented here, further trace element and Hf–Nd isotope data are reported by Paulick et al. (2010). The trace element compositions of the MAR South samples nearly cover the whole compositional range of global MORB, tapping variably depleted mantle domains (Paulick et al., 2010). Other MORB samples analysed in this study comprise glasses from the Pacific–Antarctic Rise (Castillo et al., 1998), the East

The Earth’s tungsten budget during mantle melting and crust formation

Pacific Rise (Niu and Batiza, 1997; Regelous et al., 1999; Niu et al., 2002), the Southwest Indian Ridge (Haase et al., 1996), the Mohns and Kolbeinsey Ridges (Devey et al., 1994) and the Garrett Fracture Zone (Wendt et al., 1999). For the arc lavas discussed in this study, major, trace elements and Sr, Nd, Hf, Pb isotope as well as high-precision isotope dilution data for high field strength elements (including W) are provided by Ko¨nig et al. (2008) for Cyprus and the Solomon Islands and by Ko¨nig et al. (2010) for Papua New Guinea (PNG). For the Kamchatka arc suite major, trace elements and Sr–Nd–Hf–Pb isotope data are published by Dorendorf et al. (2000), Churikova et al. (2001) and Mu¨nker et al. (2004). Further data are reported for the Tonga-Kermadec arc which shows a southward increasing flux of slab components derived from subducted pelagic sediments (shown by Regelous et al., 1997, and references therein). Further high-precision W data are presented here for the Sunda arc where large volumes of pelagic sediments are being subducted (Turner and Foden, 2001) and for the New Britain arc, which shows almost no influence of subducted sediment (Woodhead et al., 1998). All arc suites span a representative range covering settings dominated by slab fluid- and slab melt-like subduction components as indicated by different ratios of Ba–Th, Sb– Ce, Sr–Y, and Gd–Yb, respectively (e.g., Mu¨nker et al., 2004; Ko¨nig et al., 2010, and references therein). The arc suites studied cover a wide range of subducted pelagic sediment fluxes, as evident from of 207Pb/204Pb isotope variations. In addition to W, concentrations of Ta were acquired for MORB and OIB samples during the course of this study. This approach allows to test for potential sample heterogeneity by comparison to previously published high precision Ta data for these rocks (Bu¨chl et al., 2002; Mu¨nker et al., 2003; Pfa¨nder et al., 2007). In particular for small sample sizes (<100 mg), significant variations in concentrations of W and Ta may occur due to possible sample heterogeneity (e.g., nugget effects; Ko¨nig et al., 2008). Therefore, all isotope dilution measurements of W and Ta concentrations were performed on the same split of P100 mg sample powder or hand-picked glasses. The full analytical procedure for HFSE concentration measurements is described by Mu¨nker et al. (2001), Weyer et al. (2002), and Ko¨nig et al. (2008). All samples were spiked using newly prepared mixed 183W–180Ta–180Hf-tracers that were calibrated against >99.9% pure AMES metals. Unless stated otherwise, all measurements were conducted using the Thermo Finnigan Neptune MC–ICP–MS at Universita¨t Bonn. External precision and accuracy as determined by multiple digestions of different rock matrices (Weyer et al., 2002; Ko¨nig et al., 2008) are typically better than ±0.5% for Ta/W (2r r.s.d.). The external reproducibility for Ta/W was obtained for homogenous glass standards and may increase due to sample heterogeneity effects (e.g., Ko¨nig et al., 2008). Samples with very low W concentrations (<10 ppb) sometimes yielded lower external reproducibilities (e.g., ±3% for 3 individual digestions of BIR-1G, Table 1). Total blanks were generally <120 pg (Ta), <180 pg (Hf), and 19–100 pg (W). Sample blank ratios for W were always better than 100 and propagated errors of

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Table 1 Reproducibilities of standard BIR-1 and glass sample MAR-7.

BIR-1 (2) (3) Mean 2 RSD in % MAR-7 (2) (30) Mean 2 RSD in %

ppm W

ppm Ta

Ta/W

0.00718 0.00739 0.00711 0.00723 3.3 0.0846 0.0840 0.0847 0.0844 0.7

0.036 0.035 0.035 0.035 2.7 0.480 0.487 0.471 0.479 2.7

5.0 4.8 4.9 4.9 3.3 5.7 5.8 5.6 5.7 2.9

3  100 mg of each sample material was individually digested and 10% aliquot taken for each W analysis.

blank uncertainties were always better than 1%. Concentrations of Ba, Th, U, Sb, Ce were analysed using an Agilent 7500cs quadrupole ICP–MS at Universita¨t Kiel. For quadrupole ICP–MS measurements, ca. 100 mg of sample were digested and analysed in a 2% HNO3 solution. Prior to analysis, an internal In–Re standard was added. For collisional thermalization, a gas flow of 1.1 He/min was applied in the reaction cell, leading to an improved internal precision of <6% r.s.d. (2r) for Sb and Ba, and <2% r.s.d. (2r) for U, Th and Ce. This precision was regarded as sufficient to resolve the observed variations, and therefore no further isotope dilution measurements were conducted for these elements. A comparison of results for the BIR standard with recommended values of Govindaraju 1994 is given in Ko¨nig et al. (2007). Garbe-Scho¨nberg (1993) provide additional details regarding the analytical techniques. 3. RESULTS Data for W, Ta, Ba, Th, and U concentrations in MORB, OIB, and arc lavas (n = 175) are reported in Tables 2–4. Tungsten concentrations of the OIB, MORB and arc lavas range from 0.17 to 2.0 ppm (OIB), from 0.005 to 0.70 ppm (MORB), and from 0.010 to 0.18 ppm (arc lavas), respectively. These W concentration ranges broadly overlap with previously reported values for OIBs by Takamasa et al. (2009) and for OIBs, MORBs and arc rocks by Arevalo and McDonough (2008). There is no difference in W systematics between HIMU-, FOZO- and EM2-type OIBs. Ratios of W/Th in both OIB and MORB samples broadly overlap with values ranging from 0.040 to 0.23 (OIB) and from 0.090 to 0.24 (MORB). Notably, W/Th in the OIB samples range towards lower values than those in MORB. Moreover, W/Th of many arc samples also overlap with OIB and MORB values, but extend to much higher W/Th of up to 1. As for W/Th, W/U of the arc rocks also extend to higher values of up to 1.9 compared to the upper limits of W/U in OIBs and MORBs (0.2 and 0.7). Ratios of Ta/ W generally range from 3 to 10 in OIBs (with one exception of 16 in a Hawaiite from Rurutu) and from 3 to 11 in MORBs. In contrast, Ta/W in arc lavas are generally confined to ratios below 3.2 (Fig. 1a and b). Whereas Hf/W of arc lavas and OIBs all remain below 50 (with the exception of one hawaiitic OIB with a Hf/W of 263), Hf/W of

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Table 2 Selected trace element data for MORB samples. Ba, Th, U analyzed by quadrupole ICP-MS. Hf, Ta, W analyzed by high-precision isotope dilution and MC-ICP-MS. Sample

Origin

ppm Ba

ppm Th

ppm U

ppm Hf

ppm Ta

ppm W

W/Th

W/U

Ta/W

MAR-2 MAR-3 MAR-4 MAR-5 MAR-6 MAR-7 MAR-8 MAR-9 MAR-10 MAR-11 MAR-13 MAR-15 MAR-16 MAR-17 MAR-18 MAR-19 MAR-25 MAR-26 MAR-27 MAR-28 MAR-29 MAR-30 MAR-31 MAR-32 MAR-33 MAR-34 MAR-35

Segment A1 Segment A1 Segment A1 Segment A1 Segment A1 Segment A1 Segment A4 Segment A4 Segment A4 Segment A2 Segment A2 Segment A2 Segment A2 Segment A3 Segment A3 Segment A3 Turtle Pits Turtle Pits Turtle Pits Turtle Pits Turtle Pits Turtle Pits Turtle Pits Turtle Pits Turtle Pits Turtle Pits Turtle Pits

42.3 18.4 116 20.2 20.0 51.2 41.3 16.3 17.4 86.3 16.4 6.81 45.3 105 96.4 10.3 3.39 6.48 5.83 7.20 12.7 11.5 2.84 15.2 5.58 5.91 5.62

0.510 0.200 1.26 0.199 0.191 0.393 0.418 0.170 0.210 1.03 0.210 0.081 0.470 1.01 1.02 0.107 0.030 0.040 0.074 0.093 0.130 0.114 0.030 0.180 0.077 0.070 0.071

0.150 0.060 0.400 0.062 0.072 0.119 0.127 0.060 0.060 0.315 0.070 0.026 0.150 n.a. 0.295 0.035 0.011 0.010 0.031 0.042 0.054 0.048 0.010 0.051 0.031 0.060 0.060

2.101 1.632 2.675 1.734 1.443 2.246 2.463 1.670 1.792 4.116 2.424 1.179 1.894 3.610 4.105 1.481 0.5990 1.665 2.270 1.873 2.009 2.074 1.659 1.559 2.583 1.421 1.430

0.395 0.158 1.11 0.238 0.190 0.479 0.414 0.214 0.171 0.964 0.196 0.069 0.450 0.959 0.911 0.0951 0.0301 0.0310 0.0970 0.0790 n.a. 0.138 0.0310 0.141 0.100 0.0700 0.0710

0.0932 0.0350 0.220 0.0394 0.0398 0.0844 0.0803 0.0358 0.0346 0.173 0.0356 0.0145 0.0862 0.178 0.180 0.0200 0.00524 0.00579 0.00823 0.00815 0.0187 0.0163 0.00544 0.0286 0.0100 0.00924 0.00716

0.183 0.175 0.175 0.199 0.208 0.204 0.192 0.211 0.165 0.168 0.170 0.180 0.183 0.177 0.176 0.186 0.174 0.145 0.111 0.087 0.143 0.143 0.181 0.159 0.131 0.132 0.101

0.621 0.583 0.550 0.636 0.556 0.674 0.631 0.597 0.577 0.548 0.509 0.556 0.575 n.a. 0.610 0.579 0.476 0.579 0.262 0.194 0.344 0.341 0.544 0.561 0.319 0.154 0.120

4.24 4.51 5.05 6.04 4.77 5.68 5.16 5.98 4.94 5.57 5.51 4.76 5.22 5.39 5.06 4.76 5.74 5.35 11.8 9.69 n.a. 8.47 5.70 4.93 10.0 7.58 9.92

P311B PAR 14-1 PAR 631A 15 2 75-1 RR 93-7 GN 13-6 GN 13-8 GN 11-04 SWI 4-20 SWI 47-2 SWI 44-9 21853-2 23287

Pac.- Ant. Rise Pac.- Ant. Rise Pac.- Ant. Rise East Pac. Rise East Pac. Rise East Pac. Rise Garrett Frac. Z. Garrett Frac. Z. Garrett Frac. Z. SW Indian Ridge SW Indian Ridge SW Indian Ridge Kolb. Ridge Mohns Ridge

n.a. n.a. n.a. n.a. n.a. 25.1 6.99 8.60 1.16 n.a. n.a. n.a. n.a. n.a.

n.a. n.a. n.a. 2.80 1.04 0.356 0.148 0.189 0.033 n.a. n.a. n.a. n.a. n.a.

n.a. n.a. n.a. 0.927 0.408 0.141 0.072 0.290 0.178 n.a. n.a. n.a. n.a. n.a.

2.059 2.377 3.133 5.991 12.03 2.404 2.730 3.300 n.a. n.a. 1.579 n.a. 1.222 3.020

0.182 0.0960 0.306 2.36 0.799 0.588 0.170 0.182 0.0410 0.0580 0.103 0.154 0.130 0.805

0.0470 0.0338 0.0714 0.672 0.132 0.0836 0.0213 0.0260 0.00716 0.0105 0.0174 0.0180 0.0378 0.209

n.a. n.a. n.a. 0.240 0.126 n.a. 0.493 0.144 0.138 0.214 n.a. n.a. 0.235 n.a.

n.a. n.a. n.a. 0.725 0.322 n.a. 1.030 0.295 0.090 0.040 n.a. n.a. 0.593 n.a.

3.87 2.84 4.29 3.51 6.05 7.03 7.98 7.00 5.73 5.52 5.92 8.56 3.44 3.85

Ba-Th-U data for MAR samples from Paulick et al. (2010), for the Pacific-Antarctic Rise by Castillo et al. (1998), for the East Pacific Rise by Niu and Batiza (1997), Regelous et al. (1999), Niu et al. (2002), for the Southwest Indian Ridge by Haase et al. (1996), the Mohns and Kolbeinsey Ridges by Devey et al. (1994), and for the Garrett Fracture Zone by Wendt et al. (1999).

MORBs cover a much wider range from 10 to 280 (Figs. 1c and 5a). There are also clear positive correlations between Ta/W and Hf/W for MORB, with the highest Ta/W being displayed by the most depleted samples with high Hf/W. Although confined to much lower Hf/W, OIBs define a similar scatter in Ta/W as MORBs (Fig. 1c). No clear co-variation is observed between W/Th and W contents in MORB samples, whereas W/U in MORBs tend to decrease with decreasing W contents (Fig. 2a and b). Log–log diagrams of W vs. other incompatible elements (e.g., Th, U, Ba, Ta) all show slopes lower than 1 (Fig. 3). All arc rocks display well defined positive co-variations between W/Th and

W/U and negative co-variations between W/Ba and Ba (Fig. 4). Arc rocks lack a clear correlation between Ta/W vs. Hf/W, but exhibit a clear offset towards lower Ta/W if compared to OIBs and MORBs (Fig. 5a). Altogether, the above observations point to different processes controlling Ta/W, W/U and W/Th fractionations in MORBs and OIBs. Moreover, arc rocks appear to show a systematic enrichment of W relative to Ta and in some cases also relative to U and Th. The arc suites examined here comprise not only slab fluid- but also slab meltcontrolled trace element signatures (as indicated by their broad range of Ba/Th, Sb/Ce, Sr/Y, and Gd/Yb, see also

The Earth’s tungsten budget during mantle melting and crust formation

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Table 3 Selected trace element data for OIB samples. Ba, Th, U analyzed by quadrupole ICP–MS. Hf, Ta, W analyzed by high-precision isotope dilution and MC–ICP–MS. ppm Ba

ppm Th

ppm U

ppm Hf

ppm Ta

ppm W

W/Th

W/U

Ta/W

Rurutu old (HIMU) 74-386 Basalt 74-390 Basalt RR 01 Basalt RR 03 Hawaiite

Rock type

n.a. n.a. 162 256

2.35 4.49 3.15 3.74

0.70 1.02 0.87 1.01

4.230 5.196 4.922 49.60

1.80 2.60 2.86 3.08

0.311 0.373 0.425 0.165

0.13 0.08 0.13 0.04

0.44 0.37 0.49 0.16

5.79 6.97 6.73 18.7

Rurutu young (HIMU) RRT 60 Hawaiite 74-392 Hawaiite 120D Hawaiite

406 n.a. 440

6.70 5.83 7.54

1.83 1.84 2.07

7.805 8.131 9.006

4.63 4.82 5.24

0.841 0.927 1.77

0.13 0.16 0.23

0.46 0.50 0.86

5.51 5.20 2.96

Tubuai (HIMU) K109 Basanite 5433 Basanite 5434 Basalt TBA 36 Basalt 110B Nephelinite

359 440 250 226 321

4.35 8.39 4.08 3.51 19.5

0.98 2.28 1.07 0.97 4.81

5.315 6.707 4.953 4.013 9.674

4.47 5.68 3.18 2.63 10.3

0.886 1.02 0.456 0.645 1.97

0.20 0.12 0.11 0.18 0.10

0.90 0.45 0.43 0.66 0.41

5.05 5.57 6.97 4.08 5.23

Samoa (EM2) U23P Basanite U35M Basanite U13F Basalt U24L Basalt U39F Basalt U43F Basalt S46F Hawaiite

677 429 259 486 305 213 224

6.73 3.81 3.17 7.17 5.08 2.80 2.41

1.43 0.84 0.76 0.93 1.26 0.82 0.57

6.529 6.175 4.501 5.611 8.384 7.731 4.096

4.04 2.94 2.36 2.94 3.46 2.49 1.69

1.09 0.542 0.370 0.367 0.409 0.421 0.225

0.16 0.14 0.12 0.05 0.08 0.15 0.09

0.76 0.65 0.49 0.39 0.32 0.51 0.39

3.71 5.42 6.38 8.01 8.46 5.91 7.51

Society Islands (EM2) Mu5 Basalt MU 20 Hawaiite MU 29 – RI66 –

388 462 n.a. n.a.

5.13 4.80 n.a. n.a.

1.34 0.88 n.a. n.a.

6.744 11.30 6.677 8.096

2.14 3.67 2.32 2.21

0.326 0.365 0.726 0.522

0.06 0.08 n.a. n.a.

0.24 0.41 n.a. n.a.

6.56 10.1 3.20 4.23

Ba, Th, U data for Rurutu by Chauvel et al. (1995), for Tubuai by Chauvel et al. (1992), for Samoa by Workman et al. (2004), and for the Society Islands by Blais et al. (2002). Hf data by Pfa¨nder et al. (2007).

Defant and Drummond, 1990; Jochum and Hofmann, 1997; Mu¨nker et al., 2004; Ko¨nig et al., 2010). Elevated Ba/Th at low Th abundances and high Sb/Ce indicate a predominant influence of slab fluids, whereas high Sr/Y and Gd/Yb are explained by the presence of slab melts with co-existing residual garnet. Notably, the selective enrichment of W relative to both Th and U is coupled to increasing Sb/Ce (Fig. 4) and appears to be related to slab fluidcontrolled arc regimes as previously suggested by Ko¨nig et al. (2008). In contrast, arc lavas with low W/Th and W/U also show low Sb/Ce and these lavas are most likely dominated by slab melts in their sources. In W/Th vs. Ta/W space (Fig. 5b), both MORBs and OIBs define negative co-variations with the OIB samples defining a much steeper array. The OIB samples follow a trend of selective W loss. The slab fluid-related arc lavas that exhibit elevated W/Th ratios may be seen as a continuation of the OIB array. Following previous models for OIB sources (e.g., Hofmann and White, 1982), both groups might represent complementary endmembers where OIBs are depleted in fluid mobile elements (such as W) and arc lavas are enriched. The primitive mantle value is likely to lie between these two endmember groups.

4. DISCUSSION 4.1. Behaviour of W during mantle melting Tungsten has long been regarded as a highly incompatible lithophile trace element during mantle melting and silicate differentiation, quite similar to Ba, Th, or U (e.g., Rammensee and Wa¨nke, 1977; Newsom and Drake, 1979; Palme and Rammensee, 1981). Such a high incompatibility of W during mantle melting has recently been questioned by Babechuk et al. (2010). Based on a study of ophiolitic perodotites from Permian massifs in the North American Cordillera (Canil et al., 2006), Babechuk et al. (2010) proposed that W may behave much more compatible due to the presence of residual W-rich alloys in the Earth’s mantle. However, as the W enrichments in the ophiolitic peridotites are coupled with significant enrichments of LILE (Fig. 5 of Babechuk et al., 2010) it appears more than plausible that the samples have been overprinted by percolating melts or fluids. Therefore, the peridotite compositions may not represent depleted mantle residues left after melt extraction. In addition to this observation, in Ta/W vs. Hf/W space (Fig. 5a) the compositions of many terrestrial basalts do

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Table 4 Selected trace element data for arc samples. Ba, Th, U, Sb, Ce analyzed by quadrupole ICP-MS. Hf, Ta, W analyzed by high-precision isotope dilution and MC-ICP-MS. Island

Rock type

ppm Ba

ppm Th

ppm U

ppm Sb

ppm Ce

ppm Hf

ppm Ta

ppm W

W/Th

W/U

Ta/W

67 70 75 77a 78 79 82 93 96 98 99 100 102 104 106 112 89 113 115 119 122 124

Goodenough Goodenough Goodenough Goodenough Goodenough Goodenough Amphlett-Lawa Am-Watota Am-Dilia Am-Yabwaia Am-Yabwaia Am-Noapoi Am-Tuboa Am.Wamea Am-Wata Am Tewara Fergusson Fer-Sanaroa Fergusson Fergusson Normanby Normanby

Andesite Basalt bas. And. Basalt bas. And. bas. And. Rhyolite bas. And. Andesite Dacite Andesite Andesite Andesite Andesite bas. And. bas. And. Rhyolite Rhyolite Rhyolite Rhyolite Basalt Rhyolite

560 409 559 483 767 592 849 640 613 670 629 1059 1132 928 685 712 900 796 65.6 7.74 194 550

5.14 2.38 5.88 3.64 7.1 3.83 8.45 4.6 5.78 5.8 5.42 13.5 14.1 11.1 5.27 6.26 8.16 10.2 10.8 15.5 1.4 10.3

1.29 0.371 1.51 0.905 2.55 0.846 2.06 1 1.5 1.66 1.42 3.86 3.7 3.07 1.33 1.5 1.98 3.12 3.69 5.9 0.463 2.95

0.119 0.0834 0.0925 0.0598 0.292 0.0453 0.122 0.0972 0.136 0.131 0.166 0.277 0.304 0.253 0.0942 0.151 0.101 0.208 0.332 0.439 0.0651 0.17

37.7 46.8 57.2 51.9 66.9 60.9 29.1 64.1 39.8 41.3 43.4 70.7 93.3 70.1 60.5 63.9 27.9 63.3 61.9 88.1 32.9 29.9

3.665 3.293 4.396 3.705 5.416 3.419 7.391 5.079 3.529 3.708 3.983 4.525 4.87 4.8 4.662 4.926 6.541 9.151 20.35 32.48 4.025 7.406

0.257 0.301 0.421 0.304 0.564 0.172 0.655 0.472 0.278 0.394 0.455 0.403 0.401 0.466 0.527 0.518 0.582 0.817 1.53 2.21 0.379 0.899

0.495 0.143 0.471 0.238 1.05 0.188 0.86 0.408 0.345 0.4 0.357 0.618 0.729 0.623 0.439 0.49 0.783 1.24 1.24 1.68 0.117 1.12

0.096 0.06 0.08 0.065 0.147 0.049 0.102 0.089 0.06 0.069 0.066 0.046 0.052 0.056 0.083 0.078 0.096 0.121 0.115 0.108 0.083 0.108

0.384 0.385 0.311 0.262 0.41 0.222 0.418 0.408 0.23 0.241 0.252 0.16 0.197 0.203 0.331 0.326 0.396 0.397 0.337 0.284 0.253 0.38

0.519 2.105 0.895 1.281 0.54 0.918 0.761 1.157 0.806 0.985 1.276 0.652 0.55 0.747 1.2 1.057 0.744 0.661 1.231 1.321 3.239 0.803

New Britain SHR8 SHR9 SHR12 SHR 31 SHR24 SHR26 SHR14 SHR15 SHR16 SHR18 SHR21

Volcano Tavuvur Tavuvur Tavuvur Tavuvur Sulphur creek Sulphur creek Talvat Talvat Talvat Talvat Watom

Rock type Andesite Andesite Andesite bas. And. bas. And. Andesite Basalt Basalt Basalt Basalt bas. And.

ppm Ba 319 307 330 376 360 259 87.1 105 125 108 209

ppm Th 1.54 1.43 1.56 1.82 1.71 1.16 0.255 0.263 0.399 0.336 0.929

ppm U 0.847 0.793 0.864 0.98 0.936 0.62 0.14 0.136 0.212 0.155 0.518

ppm Sb 0.285 0.236 0.262 0.281 0.27 0.156 0.0625 0.0549 0.0738 0.0564 0.153

ppm Ce 25 24.7 26.5 30.1 28.6 23 8.57 10.1 11.5 9.18 15.2

ppm Hf 3.44 2.96 3.18 3.55 3.43 2.86 0.862 0.93 1.22 0.969 1.62

ppm Ta 0.163 0.139 0.148 0.167 0.161 0.118 0.042 0.045 0.053 0.045 0.071

ppm W 0.48 0.771 0.459 0.477 0.492 0.348 0.154 0.117 0.151 0.141 0.266

W/Th 0.312 0.541 0.293 0.263 0.287 0.299 0.605 0.445 0.378 0.42 0.287

W/U 0.567 0.972 0.531 0.486 0.526 0.561 1.1 0.861 0.711 0.912 0.514

Ta/W 0.339 0.18 0.323 0.35 0.326 0.341 0.27 0.382 0.349 0.316 0.266

Sunda Arc I2Kr2 I3Ga1 I4Sm1 I6Bu1 I9Me2 I10Wa1 I20Ke1

Volcano Krakatau Galunggung Semeru Buntu Merapi Watu Burik Kelut

Rock type bas. And. Basalt Basalt bas. And. Basalt Basalt bas. And.

ppm Ba 181 88.8 297 410 465 610 471

ppm Th 3.08 0.817 5.69 6.15 6.47 6.56 0.943

ppm U 0.705 0.192 1.34 1.37 1.31 1.08 0.366

ppm Sb 0.0872 0.103 0.201 0.283 0.168 0.234 0.158

ppm Ce 33.8 9.85 35.4 36.7 35.0 30.0 12.5

ppm Hf 3.22 1.30 3.13 3.73 2.29 1.95 1.35

ppm Ta 0.193 0.106 0.465 0.309 0.206 0.143 0.0721

ppm W 0.281 0.209 0.730 0.871 0.644 0.719 0.301

W/Th 0.0911 0.256 0.128 0.142 0.100 0.110 0.320

W/U 0.398 1.09 0.543 0.636 0.491 0.667 0.824

Ta/W 0.689 0.507 0.637 0.355 0.319 0.199 0.239

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D‘Entrecasteaux

Ta/W 0.992 0.938 0.662 1.06 0.74 0.136 W/U 0.251 0.238 0.268 0.24 0.243 1.94 W/Th 0.156 0.132 0.148 0.118 0.173 1.36 ppm W 0.124 0.093 0.0863 0.186 0.106 0.713 ppm Ta 0.123 0.0872 0.0571 0.197 0.0784 0.0968 ppm Hf 2.67 1.94 1.46 2.36 1.57 2.07 ppm Ce 21.1 17 13.1 16.6 13.5 16.6 ppm Sb n.a. n.a. n.a. n.a. n.a. n.a. ppm U 0.494 0.391 0.322 0.774 0.437 0.367 ppm Th 0.796 0.705 0.584 1.57 0.613 0.524 ppm Ba 434 307 327 419 241 307 Arc segment CKD CKD EVF EVF EVF CKD Kamchatka KLU-12 KLU-15 GAM-28 KIZ-24 SHM-04 2330

Rock type bas. And. Basalt Basalt bas. And. bas. And. Basalt

Ta/W 0.254 0.491 0.385 0.211 0.17 W/U 0.396 0.804 0.382 0.977 0.296 W/Th 0.342 0.305 0.304 0.416 0.171 ppm W 0.0551 0.0933 0.0447 0.0977 0.299 ppm Ta 0.014 0.0458 0.0172 0.0206 0.0507 ppm Hf n.a. n.a. n.a. n.a. n.a. ppm Ce n.a. n.a. n.a. n.a. n.a. ppm Sb n.a. n.a. n.a. n.a. n.a. ppm U 0.139 0.116 0.117 0.1 1.01 ppm Th 0.161 0.306 0.147 0.235 1.75 ppm Ba n.a. n.a. n.a. n.a. n.a. Arc segment Central Tonga North Tonga Central Tonga Kermadec Kermadec Tonga-Kermadec TOF-52 NTT-18 LAT-26 K37-3 K91-4

Rock type Basalt Basalt Basalt Basalt bas. And.

Baluron Agung Rindjani Muriah Ringgit I26Bl1 I28Ag1 I32Ri3 I12Mu2 I25Pe1

Basalt bas. And. Basalt Absarokite Absarokite

271 230 337 743 2151

1.39 2.70 1.78 10.6 14.0

0.360 0.652 0.454 2.14 3.64

0.0539 0.0644 0.0574 0.134 0.156

23.8 28.6 22.5 138 133

2.33 2.75 1.76 5.14 3.14

0.208 0.230 0.162 2.07 1.35

0.166 0.284 0.181 1.07 2.09

0.120 0.105 0.102 0.100 0.149

0.462 0.435 0.399 0.498 0.574

1.25 0.811 0.895 1.94 0.646

The Earth’s tungsten budget during mantle melting and crust formation

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not plot along a trend predicted for selective retention of W during mantle melting. In this case, the variation of Ta/W ratios in MORB, for example, would be of similar magnitude as the variation in Hf/W which is not observed. Hence, there is little evidence from our data for a more compatible behaviour of W during mantle melting resulting from residual metal phases in the mantle as suggested by Babechuk et al. (2010). Despite the general consensus regarding the highly incompatible behaviour of W, the relative compatibilities of W, Th, U, and Ba are still controversial. The pioneering work by Newsom et al. (1996) has presented W–Th–U data (obtained via instrumental neutron activation analysis with reproducibilities between ±10 and ±30%, 2r) on a limited set of oceanic basalts (n = 20) and altered mantle nodules (n = 7), suggesting broadly similar compatibilities of W, Ba, Th and U. In a more recent Laser-ICPMS study, Arevalo and McDonough (2008) investigated a larger suite of samples (n = 86), comprising MORBs, OIBs, and arc rocks. They reported “relatively constant” W/Th, W/Ba and W/U for a range of W concentrations in their samples, independent of the tectonic setting. Using log–log co-variation diagrams, with concentrations of Ba, Th, and U in MORB samples plotted versus W concentrations, Arevalo and McDonough (2008) show that W and U yield slopes that are closest to unity, whereas Th and Ba have slopes greater than unity. These patterns suggest that W is similarly compatible than U, whereas it is more compatible than Ba and Th. Hence, U would then represent the best geochemical analogue to W. Although the authors assert that W/U does not appear to be affected by magmatic differentiation and mantle source metasomatism, they report outliers of unusually high W/U and W/Th. These may be ascribed to sample heterogeneity effects, as the measurements were done by Laser-ICPMS where only small volumes of sample were analysed. However, instrumental, rather than sample heterogeneity issues explaining the larger scatter in the LA–ICP–MS data cannot be ruled out, although this currently seems to lack clear evidence (Arevalo, personal communication). In contrast to the above findings claiming near constant W/U and W/Th in terrestrial silicate reservoirs, our highprecision W data reveal systematic differences in W concentrations relative to those of Th, U, Ta, and Ba. We observe significant variations of Ta/W, W/U as well as W/Th between our investigated sample suites that reflect the different tectonic settings (Figs. 1, 3 and 4). Arc magmas consistently show lower Ta/W and extend to higher W/U and W/Th than OIBs and MORBs. Moreover, the W/U and W/Th of many MORBs are confined to relatively low values and both ratios display a slightly increase with W concentrations (Fig. 2a and b). These patterns indicate that W is slightly more incompatible than Th and U, in contrast to previous suggestions (e.g., Arevalo and McDonough, 2008). The incompatibility patterns are further illustrated in a log–log diagram of U, Ba, and Th vs. W. The slopes defined for our MORB samples (Fig. 3a) indicate that W is the most incompatible element among these trace elements. Barium and Th are the two elements showing nearly similar incompatibilities, whereas U and Ta are clearly less

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0.6

0.2

0

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15

s al et M

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ion at eg r eg

Fig. 2. W/U (a) and W/Th (b) versus the W concentration of all analysed MORB samples (black squares) in comparison to previous results (grey squares, Arevalo and McDonough, 2008; grey diamonds, Babechuk et al., 2010). W/Th and W/U of our MORB samples are confined to lower values compared to previous data.

nd tre

rolled Cpx cont etion trend pl source de

5 Chondrites

0 0

50

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150

200

250

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incompatible. Similar to the approach by Arevalo and McDonough (2008), OIB-type samples are evaluated separately for various reasons. First, OIBs represent a different mantle source than MORB and additional (recycled) components are potentially involved, biasing element abundances between both groups of rocks. Second, a log–log plot of OIB data alone (Fig. 3b) confirms the relative compatibilities seen in MORB data only (Fig. 3a), but with a much clearer offset between W and the other elements. Notably there is an offset between the OIB and MORB trends to elevated concentrations in OIBs. Evaluating both suites in the same plot therefore yield much higher slopes than in the original suites, unrelated to partial melting

3.0

MORB

2.0

Ba = 0.95 x W + 2.6

1.0

Th = 0.92 x W + 0.65

0.0

Ta = 0.85 x W + 0.52

-1.0 -2.0

U = 0.79 x W + 0.08

-3.0 -2.5

-2

-1.5

-1

-0.5

0

log W

(b) log Ba, Ta, Th, U

Fig. 1. Selected trace element variations illustrating the different fractionation of ratios of W relative to other incompatible lithophile elements between OIB, MORB and arcs, respectively. (a) Ta/W vs. ppm Ta, (b) Ta/W vs. ppm W. BCC = bulk continental crust. Note that only W concentrations in arc lavas overlap with OIBs, owing to the selective mobilization of W relative to Ta in subduction zones. (c) Ta/W vs. Hf/W. Incremental batch melting of a primitive mantle yields a source depletion trend with increasing Ta/W and Hf/W. The elevated ratios observed in MORB samples are produced by re-melting of this increasingly depleted source. Partition coefficients for Ta and W are from McDade et al. (2003) and for Hf from Hill et al. (2000).

log Ba, Ta, Th, U

Hf/W

3.0

OIB

2.0

Ba = 0.33 x W + 2.6 Th = 0.56 x W + 0.83

1.0

Ta = 0.53 x + 0.6488

0.0 U = 0.60 x W + 0.23

-1.0 -1

-0.5

0

0.5

log W Fig. 3. Log–log plots of Ba, Ta, Th, and U versus W for MORBs (c) and OIBs (d). A slope <1.00 indicates a higher incompatibility of W. Errors on slopes are less than 0.05 for MORB and 0.4 for OIB samples (2 sigma, respectively), thereby confirming that values for slopes are distinguishable from 1, with the exception of Ba vs W for MORB samples.

The Earth’s tungsten budget during mantle melting and crust formation

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processes. Similar to the log–log approach used by Arevalo and McDonough (2008) we note that errors on slopes are less than 0.05 for MORB and 0.4 for OIB samples (2 sigma, respectively), thereby not affecting the statistical significance of the observed relative compatibilities. The discrepancy between previously published data and our findings may be explained by the different analytical techniques employed and sample selection. The pioneering studies of Newsom et al. (1996) used instrumental neutron activation, and reproducibilities were as good as ±10%. Our isotope dilution technique permits W measurements at a typical precision and accuracy of ±1%, provided that more than 100 mg of sample are digested for analysis (Ko¨nig et al., 2008). If, however, the amount of sample consumed for an analysis is significantly lower, erratic outliers may occur even in homogenous materials such as volcanic glasses. These outliers may reflect inhomogeneous distribution of W, as already described by Palme and Rammensee

Fig. 5. Ta/W vs. Hf/W (a) and W/Th vs. Ta/W (b) variation diagrams illustrating the complementary geochemical signatures in Ta/W of OIBs (white circles) and arc rocks (white and grey triangles). Ta/W values of arc rocks are in the range of subducted sediments (SED, after Nebel et al., 2010). Archean granitoids, including TTGs (from Mohan et al. (2008)) exhibit both low Ta/W and W/Th. Note that arc lavas dominated by melt-like slab components in their sources (grey triangles) tend to exhibit W/Th similar to MORBs or slightly lower, indicating a nearly similar compatibilities of both elements. In contrast, fluid-related arc rocks (white triangles) exhibit W/Th of up to 1. The melt depletion trend defined by MORBs (black squares) is controlled by cpx melting, consistent with experimental data. OIBs follow a different trend, indicating selective W depletion, most likely inherited from dehydration during subduction. The big grey square is the new proposed depleted upper mantle field. PRIMA from Palme and O’Neill (2003).

(1981) and confirmed by Ko¨nig et al. (2008) and the quantities of analysed material are therefore crucial for optimizing accuracy and precision. The results of Arevalo and McDonough (2008) were obtained by Laser ablation ICPMS on dredged glasses, and it is possible that the smaller sample volumes ablated might lead to a larger spread in measured W concentrations. This is suggested by the somewhat larger variation in W/U and W/Th obtained for MORB by these authors (Fig. 2a and b). Again, instrumental problems during the LA–ICP–MS measurements cannot be ruled out. In analogy to our study, recently reported W–U–Th data obtained via solution quadrupole ICPMS measurements on a small suite of NE-Pacific MORB (n = 11; Babechuk et al., 2010; Fig. 2a and b) are also confined to a smaller range of W/U and W/Th. To assess the effects of partial melting on W–Th–U–Ta systematics in basalts further, a sample suite from the MAR from 7°30’S to 11°30’S (MAR South) which nearly covers the whole compositional spread of global MORB (e.g., in La/Yb) has been studied in more detail. Trace element and isotope evidence (Paulick et al., 2010) indicate that the observed compositional spread is caused by melting

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of variably depleted mantle domains. As for depletion parameters such as Zr/Nb (5.7–115) and Hf/W (>100), the Ta/W vs. Hf/W array for the MAR South samples overlaps the trend defined by global MORB data (Fig. 1c). Notably, the effects of variable degrees of melting are negligible for highly incompatible element ratios. Thus, different degrees of partial melting cannot explain the large differences in Ta/W or W/Th between MORB/OIB and arc rocks alone (e.g., Hofmann, 2003, and references therein). This picture is different for Hf/W where both elements show markedly different compatibilities during mantle melting (Hill et al., 2000; McDade et al., 2003). A mantle depletion curve modelled by extraction of melts produced via melting of variably depleted mantle reservoirs (Fig. 1c), however, can show that melts from variably depleted mantle sources can readily account for the observed Ta/W vs. Hf/W trend of our MORB suite. Based on these relationships, we conclude that variable degrees of upper mantle depletion are the key parameters to control potential variations of W relative to other highly incompatible lithophile elements. Collectively, the slight variations of W/Th, W/U and Ta/W within the MORB samples therefore result from the higher incompatibility of W relative to that of Th, U and Ta during mantle depletion. Our inferences made for W, Th, U, Ba, and Ta are consistent with experimental data (e.g., Hill et al., 2000; McDade et al., 2003), broadly suggesting a compatibility order of W < Th–U < Ta. The experimental data furthermore suggest that among the major minerals in the asthenospheric mantle, clinopyroxene is most likely to control any potential fractionation of W, Th, U, Ba, and Ta, as it is the most important carrier of these elements. In typical mantle clinopyroxene, Ta and W are both highly incompatible (D < 0.1, Hill et al., 2000; McDade et al., 2003), but Hf is only moderately incompatible (D ca. 0.5, McDade et al., 2003). In the presence of clinopyroxene, Hf/W may therefore be significantly fractionated in co-existing melts. Although DW/DTa in mantle clinopyroxene is lower than unity (ca. 0.09, Hill et al., 2000), the absolute D values (0.005 and 0.06, respectively, Hill et al., 2000) are too low to account for a significant Ta/W fractionation in co-existing melts. This prediction is clearly confirmed by the rather small Ta/W variations defined by terrestrial MORBs (Fig. 1c) compared to more variable Hf/W ratios. Altogether, the near constant Ta/W in MORB are therefore consistent with W being highly incompatible, similar to Th, U or La (Palme and Rammensee, 1981; Newsom et al., 1996; Mu¨nker et al., 2007; Arevalo and McDonough, 2008), but are inconsistent with a globally important role of residual W-rich alloys or metals (Babechuk et al., 2010) that would fractionate Ta/W ratios during mantle melting. Given these MORB-derived constraints for the behaviour of W during partial melting, the higher W/Th, W/U and lower Ta/W as observed in the arc samples are difficult to explain by different partial melting regimes alone. It is also known that, on average, the degrees of partial melting in the hydrous mantle wedge sources of arc magmas are higher if compared to anhydrous MORB source mantle (e.g., Tatsumi and Eggins, 1995). Hence, for their higher degrees of partial melting, arc magmas would be expected to

show slightly lower ratios of W to U, Th or Ta than MORB, in marked contrast to what is observed. The high W/Th, W/U and lower Ta/W in arc lavas would require much lower average degrees of partial melting than for MORBs. Notably, MORBs and OIBs overlap in their Ta/W and W/Th ratios, although the average degrees of partial melting for OIBs are substantially lower than those of MORBs (e.g., Hofmann, 2003, and references therein). Hence, an additional process is required to account for the enriched W abundances in arc lavas. The most likely process is the addition of components from subducted slabs, either via dehydration or melt migration. 4.2. Behaviour of W during subduction processes For surface-, groundwater and hydrothermal fluids, the general mobility of W in aqueous regimes has now been well documented, based on positive correlations between W and highly fluid-mobile elements (e.g., B, Be, and As; Kishida et al., 2004; Kamber et al., 2005; Arno´rsson and ´ skarsson, 2007). The solubility of W has been shown to O increase with temperature (Seiler et al., 2005). Moreover, the enrichment of W to ore grade concentration levels in porphyry or greisen deposits is attributed to hydrothermal fluids (e.g., Hedenquist and Lowenstern, 1994). During subduction, however, limited (if any) mobilization of W by hydrothermal fluids was documented. For instance, a generally immobile behaviour of W in subduction regimes was suggested by Noll et al. (1996), based on comparably smaller increases of W concentrations with typical fluid flux-parameters (e.g., B) in log–log variation diagrams. Moreover, the data by Noll et al. (1996) indicate that W/ Ba in arc lavas are lowered due to the preferential fluid mobility of Ba relative to W. These authors inferred that, not only relative to Ba, but also relative to U and Th, W is not significantly mobilized, and thus enriched to a similar extent in the continental crust as the three other elements. Due to its high charge-size ratio, W was finally classified as a relatively immobile HFSE in subduction systems, similar to Nb–Ta or Zr–Hf. A closer inspection of the Noll et al. (1996) database reveals that W/Th in arc rocks extend towards significantly higher values than MORBs and OIBs (up to 0.7 in arc rocks and up to 0.2 in MORBs/OIBs). The authors argued at the time that W is not significantly enriched in arc magmas relative to Th due to the apparent overlap of W/Th with typical OIB and MORB values. However, although B–Pb–As– Sb/Th ratios in arc lavas extend to slightly higher values, these ratios also overlap with MORB and OIB, but nevertheless Noll et al. (1996) still classified B–Pb–As–Sb as being mobile in subduction zones. In contrast to the interpretation of Noll et al. (1996), but not entirely in contradiction to their database, we have previously argued for a significant mobility of W in arc regimes, based on high precision W measurements in representative subduction-related magmas from the Solomon Islands and Cyprus (Ko¨nig et al., 2008). These previous findings are now supported by new high-precision data for subductionrelated suites from six additional settings (New Britain, Sunda arc, Kamchatka, Tonga-Kermadec, Papua New

The Earth’s tungsten budget during mantle melting and crust formation

Guinea, D’Entrecasteaux Islands). The new dataset for arc suites now comprises both magmas from slab melt and slab fluid dominated regimes, with different contributions from subducted pelagic sediments. As the arc suites analysed cover different endmembers of subduction-related volcanism, they are particularly suited to assess the behaviour of W at variable thermal regimes along subducting plates. Despite their different origin, all arc suites studied are lower in their Ta/ W (0.14–3.14) relative to upper mantle or MORB values (3.43–11.8; one exception is a MOR basalt from the Pacific–Antarctic Rise with Ta/W of 2.84). Only few arc samples overlap with the primitive mantle Ta/W value of 2.5 (Palme and O’Neill, 2003; Fig. 5b). A lack of correlation between Ta/W, W/U and W/Th in arc lavas with W abundances (Fig. 1a and b and Ko¨nig et al., 2008), SiO2, MgO or Zr/Nb (not shown) indicates that these element ratios are neither affected by crystal fractionation nor by partial melting. As the Ta concentrations of the analysed arc rocks overlap with MORB values it can be inferred that Ta behaves rather immobile whereas W was selectively enriched by subduction components. This observation clearly demonstrates that W cannot be classified as an immobile HFSE like Nb or Ta. Similar to the study of Noll et al. (1996), the measured W/Th of our arc suites extend towards higher values compared to MORB (Fig. 4a). A similar enrichment trend is evident for W/U. Both, W/U and W/Th are correlated (Fig. 4a), indicating that the variations are controlled by W enrichment. Moreover, co-variations of W/Th and W/ U with fluid flux parameters such as Ce/Pb (Chauvel et al., 1995) or Sb/Ce (Jochum and Hofmann, 1997) are observed in samples derived from magma sources predominantly controlled by subduction fluids (New Britain, Solomon Islands, Cyprus, Tonga, and Kamchatka). This confirms an effective mobilization of W by subduction fluids (Fig. 4b). In contrast to Th and U, measured W/Ba of our samples extend to slightly lower values compared to MORB (Fig. 4c), indicating that Ba is more mobile than W. In summary, there is clear evidence that W is significantly more mobile than HFSE, slightly more mobile than Th and U, but less mobile than Ba. Based on Mo–W systematics (Ko¨nig et al., 2008, 2010), the mobility of W in subduction fluids can furthermore be ascribed to its speciation as tungstate or chloride complexes, with little dependence on redox state and precursor material of the fluids. An important feature of our arc data is that several arc lavas exhibit similar or in some cases even lower W/Th and W/U ratios than MORB, although the Ta/W are uniformly low, implying similar incompatibilities of both element pairs. The low W/Th and W/U are indeed observed for arc magmas that show no significant fluid-induced fractionation of ratios such as Sb/Ce (Fig. 4b) and that were previously related to the presence of slab melts (PNG boninites, Ko¨nig et al., 2010) and/or sediment-derived melt-like components in their source (Sunda and Kermadec arc lavas, Regelous et al., 1997; Turner and Foden, 2001). In conjunction with this scenario, Archean TTG suites that are attributed to a slab melt origin also display very low W/Th, within the range of MORB (Fig. 5b). Again, the Ta/W in TTGs exhibit similarly low values than modern arc magmas

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(Mohan et al., 2008). Notably, the low W/Th in some of these samples even require that W might be even more compatible during slab melting than Th. So far, this pattern is not confirmed by the few experimental partitioning studies available for relevant minerals, but this picture might change once more partitioning data for W become available, in particular for garnet. It is known from existing partitioning studies, however, that DTh in garnet is highly dependent on the major element composition of residual garnet (Van Westrenen et al., 1999). In summary, it becomes obvious that the fractionation behaviour of W/Th and W/U during subduction processes is ambiguous, depending on the nature of the slab components involved unlike Ta/W that are uniformly low in all arc magmas. Ratios of W/Th and W/U in arc magmas might therefore be a potential discriminant between meltlike and fluid-like subduction components in contrast to Ta/W that appears to be much more insensitive to the type of subduction processes involved. However, Ta/W is a more reliable parameter to assess the general role of subduction components during crust mantle differentiation. 4.3. Tungsten geochemistry during evolution of the Earth’s silicate reservoirs The presence of three major reservoirs in the silicate Earth – primitive mantle, depleted mantle, and continental crust is a fundamental geochemical paradigm (see Hofmann, 1997, and references therein) that can successfully explain many trace element distribution patterns in the silicate Earth. During primary crust-mantle differentiation, distinct fractionations of similarly compatible trace elements are observed. For instance Nb/U and Ce/Pb of the continental crust are lower than values for PRIMA (Hofmann, 1986; Miller et al., 1994; Chauvel et al., 1995), although each of these element pairs exhibit nearly identical compatibilities during partial melting processes in the mantle. The cause for fractionation of Nb from U and Ce from Pb is their different behaviour during subduction processes with the fluid-mobile U or Pb being more enriched in the sources of arc magmas, and ultimately in the crust. Likewise, elements like Nb and Ce are preferentially retained in the mantle (e.g., Hofmann, 1988; McCulloch and Gamble, 1991; Rudnick et al., 2000; Mu¨nker et al., 2003, and references therein). The same systematics as for Ce/ Pb and Nb/U have been shown above to be true for Ta/ W. Our Ta/W data indicate that the three major reservoirs, continental crust, depleted upper mantle and primitive mantle also exhibit distinct ranges of Ta/W ratios. This is because arc rocks and globally subducted sediments show consistently lower Ta/W than OIB and MORB (Fig. 1). As demonstrated for Ce/Pb and Nb/U (Hofmann, 1986; Miller et al., 1994), the low Ta/W for arc magmas and subducted sediments also mirror the composition of the continental crust. This is supported by the limited body of high-precision Ta and W data for continental materials showing similarly low Ta/W in ocean floor sediments from the Banda arc (Nebel et al., 2010) that is regarded as being representative for recycled continental material (Vroon et al., 1995). Other estimates of Ta/W in the upper

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continental crust range from 1.1 (Wedepohl, 1995), 0.8 (Taylor and McLennan, 1985, 1995) to 0.7 (Rudnick and Gao, 2003) and 0.6 (Hu and Gao, 2008). The mechanisms active during Ta/W fractionation can be further examined using the observed complementary relationship between OIBs and some arc magmas. In W/Th vs. Ta/W space (Fig. 5b), the OIB trend most likely reflects W loss following dehydration processes in subducting oceanic crust, as the trend is matched by the W enrichment trend displayed by arc magmas dominated by fluid-like subduction components (New Britain, Solomon Islands, Cyprus, Tonga and Kamchatka). Altogether, the bimodal distribution of W/Th between fluid- and melt-controlled subduction regimes is not visible in Ta/W, as Ta behaves much more immobile during any of the two processes considered. With respect to Ta/W, a complementary relationship between all types of arc magmas and MORB–OIB is evident. The fractionation of Ta/W during subduction-related formation of the continental crust is similar to that of Nb/U (Hofmann, 1986) and Ce/ Pb (Miller et al., 1994), indicating a selective depletion of the upper mantle in U, Pb, and W relative to more immobile elements of similar compatibility (Nb, Ce, Ta). These combined parameters can therefore be used to constrain the budget of W in the silicate Earth. 4.4. Mass balance calculations of the silicate Earth’s W budget and implications for the age of core formation Due to significant fractionation of Hf from W during mantle melting (e.g., Palme and Rammensee, 1981; Righter and Shearer, 2003), the Hf/W of the terrestrial mantle cannot be measured directly and has thus traditionally been inferred from near constant W/Th and W/U ratios in silicate rocks, in conjunction with the chondritic Hf/Th and Hf/U (e.g., Newsom et al., 1996; Kleine et al., 2002; Scho¨nberg et al., 2002; Yin et al., 2002). Given the results of this study, the primitive mantle may have a substantially different W/ Th compared to the depleted upper mantle, resulting in a different estimate of Hf/W in the primitive mantle. The same systematics also apply for Ta/W in the primitive mantle. Similar to Nb/U and Ce/Pb (Hofmann, 1986; Miller et al., 1994) and also valid for Ta/W, the W/Th and W/U ratios of both OIBs and MORBs may not mirror the composition of the bulk mantle. Indeed, average Ta/W in all our arc rocks (1.0) and subducted sediments (0.7, Rudnick and Gao, 2003; Nebel et al., 2010) are lower compared to the primitive mantle (2.5; Palme and O’Neill, 2003) and an average MORB + OIB value of 5.1 (Fig. 5b). This finding is in marked contrast to previous estimates (e.g., Newsom et al., 1996; Arevalo and McDonough, 2008). A new mass balance for W must therefore take into account the selective behaviour of W during subduction zone processes. This goal is best achieved by using Ta/W systematics, as W/Th and W/U systematics cannot unambiguously identify subduction processes as outlined above. For a new mass balance calculation, only samples with near primitive mantle like Zr/Nb (<40) were taken into account in order to calculate a mean Ta/W value for MORB and OIB, as samples with Zr/Nb > 40 have experienced an extensive degree of depletion, resulting in slightly elevated

Ta/W (Fig. 1c). An improved mass balance must therefore take into account the lower W content of the depleted upper mantle which can be derived from the average Ta/W of our MORB and OIB suites (5.1). Previous studies by Salters and Stracke (2004) and Arevalo and McDonough (2008) proposed a W abundance in the depleted mantle of 3.5 ± 2 ppb and of 3.0 ± 2.3 ppb, respectively. Parameters used for mass balance calculations are listed in Table 5. The values obtained for the Earth’s primitive and depleted mantle are listed in Table 6. Compared to previous estimates of W contents in the continental crust (1 ppm, Rudnick and Gao, 2003), a slightly lower W abundance (0.7 ppm) is now calculated using the well known Ta abundance of the continental crust (0.7 ppm, Rudnick and Gao, 2003) and our crustal Ta/W (1.0) estimated from the composition of arc rocks. This estimate is slightly higher than the Ta/W value by Rudnick and Gao (2003) obtained for crustal sediments (0.7). Altogether the different studies result in a “crustal range” of Ta/W with a subsequent range of recalculated Hf/W for the primitive mantle. If a Ta/W of 0.7, inferred from crustal sediments, is now considered as the most representative estimate for the bulk continental crust, it can be used for mass balance calculations estimating the composition of the primitive mantle following the equation below modified from Hofmann et al. (1986)   ðWCC  ðTa=WÞDM  XCC Þ þ TaPM  XCC  TaCC WPM ¼ ðTa=WÞDM ð1Þ where PM is the primitive mantle, DM the depleted (upper) mantle (residual mantle in Hofmann et al., 1986), CC the continental crust and XCC the mass fraction of continental crust. Based on the revised mass balance, we obtain a W abundance of 12 ppb for the primitive mantle and of 2.4 ppb for the depleted mantle. This results in new primitive mantle values for Ta/W of 3.4, for W/Th of 0.14, for W/U of 0.53, and for Hf/W of 25.8. The new mean W concentration for the depleted mantle is slightly lower and may be more robust compared to the estimates of Salters and Stracke (2004) and Arevalo and McDonough (2008), the new concentration for the primitive mantle is also slightly lower than the values of Palme and O’Neill (2003); 16 ppb). In particular our new Hf/W value is significantly higher compared to the previously accepted value of 18.7 (Newsom et al., 1996; Palme and O’Neill, 2003). It is noteworthy, that the resulting W/Th of 0.14 ± 0.015 for the primitive mantle is slightly lower, but still within error of our MORB value (0.17 ± 0.015). If the composition of the Earth’s crust would have been controlled by subduction fluid-dominated trace element enrichment, a significantly higher W/Th would be expected for the primitive mantle compared to the depleted mantle. However, if statistically significant, the slightly lower W/Th of the Earth’s primitive mantle may mirror an important role of subduction meltdominated processes during early crust formation. As shown by our arc data, melt-controlled subduction zone processes would result in a lower crustal W/Th and a higher W/Th in the depleted upper mantle. Again, these observations confirm our previous notion that mass balances based on W/Th ratios are rendered difficult. For the mass balance

The Earth’s tungsten budget during mantle melting and crust formation Table 5 Parameters used for mass balance calculations. Parameter

Value

Source

ppm TaCC (Ta/W)CC XCC ppm TaPM (Ta/W)DM (Hf/Ta)CI

0.7 0.7 0.0054 0.040 5.1 7.50

Rudnick and Gao, 2003 this study Rudnick, 2000 Palme and O Neill, 2003 this study Palme and O Neill, 2003

PM = primitive mantle, DM = depleted mantle, CI = chondritic value, CC = continental crust and XCC the mass fraction of continental crust.

Table 6 Values calculated for the Earth’s primitive and depleted mantle. Parameter

Value

ppb WPM ppb WDM (Ta/W)PM (W/Th)PM (W/U)PM (Hf/W)PM

12 2.4 3.4 0.14 0.53 25.8

based on Ta/W, additional uncertainties are evaluated in Fig. 6. These uncertainties may lie in the mass fraction of continental crust assumed and the estimated Ta/W of the continental crust. However, a significantly higher mass fraction of continental crust than the estimated 0.54% and a significantly different Ta/W outside current estimates for the continental crust (0.7–1.0) would be required to match the previously reported primitive mantle values for Ta/W (2.5, Palme and O’Neill, 2003) and Hf/W (18.7, Newsom et al., 1996; Palme and O’Neill, 2003). Both of these scenarios are highly unlikely (Hofmann et al., 1986; Rudnick and Gao, 2003, and references therein). The Hf/W ratio of the bulk silicate Earth is an important input parameter for applying the extinct Hf/W chronometer to date terrestrial core formation. The extinct 182 Hf–182W isotope system is the method of choice because 182 Hf has a half-life of 9 Myrs, covering timescales of the Earth’s accretion and differentiation. Moreover, both elements show a contrasting behaviour during metal segregation. The lithophile Hf is almost entirely retained in the mantle, whereas the moderately siderophile W is preferentially partitioned into the metal core. A high Hf/W ratio in the early primitive mantle would result in the accumulation of more radiogenic 182W compared to chondrites and the core (Harper and Jacobsen, 1996; Kleine et al., 2002; Scho¨nberg et al., 2002; Yin et al., 2002). Assuming that the Earth’s core formation occurred in one single event, a model age of 33 Ma after formation of the solar system has been reported using the previously accepted Hf/W of 18.7 (Kleine et al., 2002; Scho¨nberg et al., 2002; Yin et al., 2002). Given the physical constrains on Earth’s accretion, core formation was probably a continuous process that was also affected by a series of larger differentiated impactors (e.g., Wetherill, 1994; Kramers, 1998). In the latter case, the 182W excess generated early in the Earth’s his-

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tory would have been partly diluted following by further accretion of undifferentiated material and the model ages would be much younger, as young as 70 Myrs after solar system formation (e.g., Halliday, 2004; Kleine et al., 2004). Another endmember scenario assumes that the Earth accreted from differentiated material with little equilibration following collision with the proto Earth (e.g., Alle`gre et al., 2008; Rudge et al., 2010). In this case, the W isotope signature would simply be inherited from the precursor material and would be of no age significance. As a consequence of these consideration, the single stage core formation age can only be regarded as a maximum age. Our revised mass balance, taking into account the calculated Ta/W of PRIMA (3.4), results in a higher Hf/W (ca. 25.8) compared to previous estimates (18.7, Newsom et al., 1996; Palme and O’Neill, 2003). As a result of this new Hf/ W value for PRIMA, the model ages obtained for formation of the Earth’s core may change significantly (see Table 7 for equation parameters and Table 8 for results). The model age assuming a simplified single stage core formation scenario can be calculated using the following equation (Lee and Halliday, 1995): 182  182   180  W W Hf DT ¼ k1  ln  = 184 W 184 W 184 W BSE CHOND BSE 180   182   Hf Hf = 180  184 W CHOND Hf SSI

ð2Þ

where k = 7.79  108 Myr is the 182Hf decay constant 182 (Vockenhuber et al., 2004) and Hf/180Hf SSI = 9.72 ± 0.44  105 is the bulk solar system initial value determined by Burkhardt et al. (2008), 180Hf/184WBSE are the Bulk Silicate Earth values (this study and Palme and O’Neill, 2003), and 180Hf/184WCHOND the chondrite value of Kleine et al. (2004). Assuming single stage core formation and using our newly proposed Hf/W of 25.8 for PRIMA, a maximum age of now 38 Myrs after solar system formation is obtained, compared to a previously reported model age of 33 Myrs as based on a Hf/W of 18.7 (Fig. 7). Following the approach of Jacobsen (2005) the model age assuming continuous (i.e., exponential) accretion can be calculated according to the following equation: 180  182  Hf Hf e182 ðtÞ ¼ 104  182  180  k  f2r=s W CHUR Hf SSI  I ar=s ðk; tÞ

ð3Þ

where 180Hf/182WCHUR is the ratio of the chondritic uniform reservoir, k is the 182Hf decay constant, f2r/s is the Hf/W fractionation factor between the Earth’s mantle and the chondritic uniform reservoir (Jacobsen, 2005; Eq. (1)), and I ar=s ðk; tÞ describes the change in fractional mass of the mantle assuming (1) exponential accretion as originally proposed by Wetherill (1994) and (2) constant mantle/core mass ratios during differentiation. A continuous addition of undifferentiated material has some important bearing on the evolution of 182W because it slows down the ingrowth of radiogenic 182W, yielding generally younger model ages compared to single stage models. If a continuous accretion scenario with constant metal/ silicate ratios including full metal silicate equilibration and a Hf/W fractionation factor (f2r/s) of 15.6 is taken into

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PRIMA

(Ta/W)

2.5

b)

26

(Hf/W)

PRIMA

3.0

30

a)

3.5

2.0 1.5 1.0 0.004

0.008

0.012

22 18 14 10 0.004

0.016

Mass fraction continental crust in %

c

a

b

x x

2.0

c

30

a x x

25 20

0.2

0.4

0.6

0.8

1

0

1.2

0.2

0.4

0.6

e) PRIMA

4.0 3.5

(Hf/W)

(Ta/W)

PRIMA

4.5

3.0 2.5

0.028

0.036 ppm Ta

0.8

1

1.2

Ta/W CC

(Ta/W) CC

0.044

0.052

34 f) 32 30 28 26 24 22 20 0.028

0.036

0.044

ppm Ta

PM

40

g)

36 PRIMA

5.0 4.0

(Hf/W)

PRIMA

x

b

10 0

(Ta/W)

0.016

15

1.0

6.0

0.012

d)

35

x

PRIMA

3.0

40

c)

(Hf/W)

PRIMA

4.0

(Ta/W)

5.0

0.008

Mass fraction continental crust in %

3.0

0.052

PM

h)

32 28 24 20

2.0

16 1.0

3.0

5.0 (Ta/W)

7.0

9.0

11.0

DM

1.0

3.0

5.0 (Ta/W)

7.0

9.0

11.0

DM

Fig. 6. Calculated Ta/W and Hf/W (curves) in the primitive mantle for variable mass fractions of continental crust (a) and (b), Ta–W of the continental crust (c) and (d), Ta–W of the depleted mantle (e) and (f), and ppm Ta of the depleted mantle (DM) (g) and (h). Grey fields indicate mean values with 2r standard error taken for the mass balance estimate. Literature values for the mean Ta/W of the continental crust are from (a) Rudnick and Gao (2003); (b) Taylor and McLennan (1995) and (c) Wedepohl (1995). See text and appendix for modelling parameters and calculation.

account (deduced from Hf/W = 18.7) a mean accretion time of 10.7 Myrs (at 63% accretion) is calculated applying the model proposed by Jacobsen (2005). If, however, a Hf/ W of 25.8 as proposed by our study is considered a longer mean accretion time of 11.4 Myrs can be derived. These estimates can be used to calculate core formation model ages assuming that core formation ceased after 99% of the Earth was accreted. This assumption is strongly sup-

ported by the late veneer model which relies on the silicate Earth’s PGE abundances being established by addition of 1% of the Earth’s mass as primitive chondritic material after core formation was completed (e.g., Halliday, 2004; Kleine et al., 2004). Therefore the end of core formation can be deduced from the accretionary mean life: tCF ¼  lnð1  F tCF Þs

ð4Þ

The Earth’s tungsten budget during mantle melting and crust formation Table 7 Summary of Parameters used for model calculations Parameter

Value

182

Source -8

k Hf I ar=s ðk; tÞ (182Hf/180Hf)SSI (182W/184W)BSE (182W/184W)CHUR (180Hf/184W)CHUR

7.79 x 10 Myr see Jacobsen (2005) (9.72 ± 0.44) x 10-5 0.864844 (0 eW) 0.864680 (-1.9 eW) 1.32

Vockenhuber et al., 2004 equation 70 Burkhardt et al.., 2008 Schulz et al., 2010 Kleine et al., 2004 Anders & Grevesse, 1989

Table 8 Results of single stage and continuous core formation models As a function of the (Hf/W)BSE. Core formation

(Hf/W)BSE

Source

Model age

single stage

18.7 25.8 18.7 25.8

Newsom et al. (1996) this study Newsom et al. (1996) this study

33 38 51 53

continuous

Myrs Myrs Myrs Myrs

where tCF is the time of core formation, s is the accretionary mean life and F is the mass fraction of the Earth when core formation ceased (i.e., 0.99). When applying this equation a model age of 51 Myrs (corresponding to s = 10.7 and Hf/ W = 18.7) can be determined, whereas the new estimate of Hf/W = 25.8 which corresponds to s = 11.4 yields a slightly younger core formation model age of 53 Myrs. In the endmember case that the Earth accreted from differentiated material with little equilibration (Alle`gre et al., 2008; Rudge et al., 2010), the revised Hf/W from our study would have little bearing on the age estimates for core formation. Such endmember scenarios, however, are rendered unlikely for the following reasons: first, such models assume highly radiogenic 182W abundances in the asteroidal silicate material accreted (e.g., Rudge et al., 2010). Recent surveys of W

3.0 2.0

PM Ta/W = 2.5 Hf/W = 18.7

ε182W

1.0 0.0

yr 33M

-1.0 -2.0

PM Ta/W = 3.4 Hf/W = 25.8

r 38My

CC -3.0 0

5

10

15

20

25

30

Hf/W Fig. 7. Isochron diagram illustrating the effect of higher Ta/W and Hf/W of the primitive mantle on the Hf/W model age of simplified single stage core formation. With previously suggested Ta/W of PRIMA = 2.5 (Palme and O’Neill, 2003) a Hf/W of 18.7 and a Hf– W model age of 33 Myrs after solar system formation are obtained. With the new Ta/W for PRIMA (3.4) and a Hf/W of 25.8 (this study) the age is 38 Myrs. See text and appendix for modelling parameters and calculation.

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isototope compositions of meteorites (see Schulz et al., 2010 and references therein), however, reveal that most smaller asteroidal bodies underwent late thermal re-equilibration. As a consequence, their 182W abundances in the silicates are much less radiogenic and the degree of metal:silicate equilibration assumed in such endmember models, in particular throughout the Earth’s early accretion, needs to be much lower than actually predicted by physical considerations (e.g., Dahl and Stevenson, 2010; Rubie et al., 2010). Secondly, a recent trace element study of lunar rocks (Mu¨nker, 2010) revealed that the Hf/W ratio of the silicate Moon (now ca. 24.9) overlaps the Hf/W ratio proposed here for the silicate Earth. Together with the identical 182 W compositions of the silicate Earth and the Moon (Touboul et al., 2007), this coincidence strongly implies that the Moon forming giant impact might have triggered an efficient metal–silicate re-equilibration on the Earth. It is therefore likely, that the model age for single stage core formation may in fact be close to the age of the Moon forming giant impact. In summary, the higher revised Hf/W value determined in this study for the silicate Earth implies, that core formation on Earth was slightly more prologued than previously thought. The shift in age caused by the new Hf/W value, however, is model dependent. The effect of the Hf/W in the silicate Earth is linked to the degree of radiogenic ingrowth of 182W in the silicate Earth itself. In all of these cases, however, the shift in age due to the revised Hf/W ratio still lies within the large uncertainties of the individual modelling approaches employed (see also Halliday, 2004; Kleine et al., 2004; Jacobsen, 2005). 5. CONCLUSION High-precision tungsten concentration data have been obtained by isotope dilution and MC–ICPMS for a broad range of mafic igneous samples (n = 175, including companion studies), covering a variety of tectonic settings including mid ocean ridge, ocean island and island arc settings. These data provide the following constraints on the W budget of silicate reservoirs on Earth and its bearing on the Hf–W model age of the Earth’s core: (1) During partial melting, a compatibility sequence in the order Ba < W < Th < U < Nb < Ta < REE can be observed, in accord with previous work by Newsom et al. (1996). Our data is inconsistent with previous interpretations that Ba and Th are more incompatible than W (Arevalo and McDonough, 2008) and models suggesting that W may behave compatibly during mantle melting (Babechuk et al., 2010). (2) Arc magmas from all investigated settings all show significantly lower Ta/W (mean = 1.0, n = 111) than OIB and MORB (3–12, n = 64) and extend towards higher W/U and W/Th. Similarly, Ta/W of globally subducted sediment, that represent the bulk continental crust, is lower (mean = 0.7) compared to MORBs and OIBs. These patterns reflect the higher mobility of tungsten in subduction zone systems,

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particularly if fluid-like subduction components and recycled continental sediments are involved. Arc magmas originating from slab melt-dominated regimes exhibit similar or even lower W/Th than MORB, calling for a higher compatibility of W relative to Th during slab melting. As W behaves higly mobile during formation of the Earth’s crust, the W/Th and Ta/W of the Earth’s depleted upper mantle are different from the Ta/W and W/Th in the primitive mantle. Due to the contrasting behaviour of W–Th and W–U in different subduction regimes, ratios of Ta/W are best suited for mass balance calculations assessing the W budget of the silicate Earth. As for Nb/U and Ce/Pb, the Ta/W ratios of arc magmas that represent the continental crust are systematically lower compared to MORBs and OIBs. This permits a more precise mass balance calculation using crustal Ta/W between 0.7 and 1.0. The new proposed abundance of W is 2.4 ppb in the depleted upper mantle and 12 ppb in the primitive mantle. This results in new primitive mantle values for Ta/W = 3.4, for W/Th = 0.14, and for W/ U = 0.53. The new Hf/W value for the primitive mantle is 25.8, significantly higher compared to the previously published value of 18.7. The new Hf/W value calculated here for the Earth’s primitive mantle has important implications for the 182 Hf–182W dating of core formation on Earth. Using the new Hf/W of the primitive mantle, the silicate Earth’s Hf–W model ages might be younger by as much as 5 Myrs. Assuming a single stage core formation model, a maximum age of 38 Ma is now obtained for the formation of the Earth’s core. The difference is smaller or even negligible using more complex core formation models (e.g., Jacobsen, 2005; Alle`gre et al., 2008; Rudge et al., 2010).

ACKNOWLEDGEMENTS This research was supported by the DFG (German Research Foundation, projects MU 1406/6 and MU 1406/7). Dieter Garbe-Scho¨nberg and Ulrike Westernstro¨er from Universita¨t Kiel are thanked for quadrupole ICP–MS analyses. A. Luguet was always helpful with solving Neptune problems. T. Schulz and J.E. Hoffmann are thanked for discussion. Arc samples were kindly provided to C. Mu¨nker by G. Wo¨rner and T. Worthington. MORB and OIB samples were kindly provided to A. Bu¨chl and J. Pfa¨nder by C. Chauvel, C. Devey, K. Haase, D. Mertz, Y. Niu, M. Regelous, and J. Snow. Two anonymous reviews and a detailed review by A. Stracke significantly helped to improve this paper.

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